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Niu, Y.L. (2016) 'The meaning of global ocean ridge basalt major element compositions.', Journal of
petrology., 57 (11-12). pp. 2081-2104.
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JOURNAL OF
PETROLOGY
Journal of Petrology, 2016, Vol. 57, No. 11&12, 2081–2104
doi: 10.1093/petrology/egw073
Original Article
The Meaning of Global Ocean Ridge Basalt
Major Element Compositions
Yaoling Niu1,2,3*
1
Institute of Oceanology, Chinese Academy of Sciences, Qingdao, 266071, China; 2Department of Earth Sciences,
Durham University, Durham DH1 3LE, UK; 3Laboratory for Marine Geology, Qingdao National Laboratory For
Marine Science and Technology, Qingdao 266061, China
*Corresponding author. Department of Earth Sciences, Durham University, Durham DH1 3LE, UK.
Telephone: þ44-19-1334-2311. Fax: þ44-19-1334-2301. E-mail: [email protected]
Received April 12, 2016; Accepted December 16, 2016
ABSTRACT
Mid-ocean ridge basalts (MORB) are arguably the most abundant and also the simplest igneous
rocks on the Earth. A correct understanding of their petrogenesis thus sets the cornerstone of igneous petrogenesis in general and also forms the foundation for studying mantle dynamics. Because
major element compositions determine the mineralogy, phase equilibria and physical properties of
rocks and magmas, understanding global MORB major element systematics is of prime importance. The correlated large MORB major element compositional variations are well understood as
the result of cooling-dominated crustal-level processes (e.g. fractional crystallization, magma mixing, melt–rock assimilation or reaction, and other aspects of complex open-magma chamber processes), but it remains under debate what messages MORB major elements may carry about mantle
sources and processes. To reveal mantle messages, it is logical to correct MORB melts for the effects of crustal-level processes to Mg# 072 to be in equilibrium with mantle olivine of Fo90.
Such corrected MORB major element (e.g. Si72, Ti72, Al72, Fe72, Mg72, Ca72 and Na72) compositional
variations thus reflect fertile mantle compositional variation, composition-controlled mantle physical property variation (e.g. density and solidus), variation in the extent and pressure of melting,
and uncertainties associated with the correction. The correction-related uncertainties can be
removed through justified heavy averaging. Because ridge axial depth variation (0 to 6000 m
below sea level) and plate spreading rate variation (<10 to >150 mm a–1) are the two largest known
physical variables along the global ocean ridge system, possible correlations of MORB major element compositions at Mg# 072 with these two physical variables are expected to reveal intrinsic
controls on global MORB petrogenesis and ocean ridge dynamics. Indeed, global MORB major
element data averaged with respect to both ridge axial depth intervals and ridge spreading rate
intervals show significant first-order correlations. These correlations lead to the conclusion that the
ridge axial depth variation and MORB chemistry variation are two different effects of a common
cause, induced by fertile mantle compositional variation. The latter determines (1) variation in both
composition and mode of mantle mineralogy, (2) variation of mantle density, (3) variation of ridge
axial depth, (4) source-inherited MORB compositional variation, (5) density-controlled variation in
the maximum extent of mantle upwelling, (6) apparent variation in the extent of melting, and (7)
the correlated variation of MORB chemistry with ridge axial depth. These correlations also confirm
the recognition that the extent of mantle melting increases with, and is caused by, increasing plate
spreading rate. Mantle temperature variation could play a part, but its overstated role in the literature results from a basic error (1) in treating ridge axial depth variation as solely caused by mantle
temperature variation by ignoring the intrinsic control of mantle composition, (2) in treating mantle
plume-influenced ridges (e.g. Iceland) as normal ridges of plate spreading origin, and (3) in treating
seismic low velocity at great depths (>300 km) beneath these mantle plume-influenced ridges as
evidence for hot ridge mantle. There is no evidence for large mantle temperature variation beneath
ridges away from mantle plumes. The suggested conclusions of this study may continue to be
C The Author 2017. Published by Oxford University Press.
V
This is an Open Access article distributed under the terms of the Creative Commons Attribution License (http://creativecommons.org/licenses/by/4.0/), which permits
unrestricted reuse, distribution, and reproduction in any medium, provided the original work is properly cited.
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Journal of Petrology, 2016, Vol. 57, No. 11&12
debated, but they are most objective, and are most consistent with petrological, geochemical, geological and geophysical principles and observations.
Key words: mid-ocean ridges; MORB chemistry variation; MORB petrogenesis; mantle melting; fertile mantle compositional control; spreading rate control
INTRODUCTION
Mid-ocean ridge basalts (MORB) are arguably the most
abundant igneous rocks on the Earth because they cover
much of the global ocean floor and continue to form
along the 60 000 km long globe-encircling ocean ridges.
They are also arguably the simplest igneous rocks on the
Earth because they have relatively uniform chemical and
isotopic compositions and because their origin is well
understood as a straightforward consequence of plate
tectonics; that is, plate separation induced passive mantle
upwelling, decompression melting, melt extraction and
cooling-dominated crustal-level differentiation (e.g. Klein
& Langmuir, 1987; McKenzie & Bickle, 1988; Niu & Batiza,
1991a; Sinton & Detrick, 1992; Niu, 1997). Nevertheless,
MORB compositional variations do exist on all scales and
have been recognized from the early days of MORB studies (e.g. Sun et al., 1975, 1979; Bryan & Moore, 1977;
Langmuir & Hanson, 1980; Schilling et al., 1983). The correlated large major element compositional variation has
been correctly explained as resulting from low-pressure
crustal-level differentiation (O’Hara, 1968a, 1968b; Walker
et al., 1979; Stolper, 1980; Perfit & Fornari, 1983; Christie
& Sinton, 1986; Langmuir et al., 1986; Langmuir, 1989;
Sinton et al., 1991; Sinton & Detrick, 1992; Batiza & Niu,
1992; Rubin et al., 2009; O’Neill & Jenner, 2012; Coogan
& O’Hara, 2015). Variations in abundances and ratios of
incompatible elements and radiogenic isotopes have
been logically interpreted as reflecting fertile mantle
source compositional variation (e.g. Frey et al., 1974; Sun
et al., 1975, 1979; Zindler et al., 1984; Mahoney et al.,
1994; Niu et al., 1996, 1999, 2001; Castillo et al., 1998,
2000).
In addition to these repeatedly verified and generally
accepted interpretations, MORB melts, in particular their
major element compositions, are predicted also to carry
information on both mantle sources and processes as
shown by the correlated variations among major element abundances, incompatible element ratios and radiogenic isotopes (e.g. Langmuir & Hanson, 1980; Christie
& Sinton, 1986; Natland, 1989; Niu & Batiza, 1991a, 1997;
Mahoney et al., 1994; Shen & Forsyth, 1995; Niu et al.,
1996, 1999, 2001, 2002a; Lecroart et al., 1997; Rubin
et al., 2009) and by experimental petrology (e.g. Jaques
& Green, 1980; Green & Falloon, 1998, 2005, 2015;
Green, 2015). On the basis of ample observations and
many studies, it is reasonable to state that MORB major
element compositional variation must be inherited from
the fertile mantle source compositional variation while
also recording the physical conditions of magma generation as a result of plate spreading rate variation (Niu &
Hékinian, 1997a, 1997b) and possibly mantle potential
temperature variation (see Niu et al., 2001). In this regard, the work by Klein & Langmuir (1987), after Dick
et al. (1984), offered an unprecedented insight. It showed
correlated variations of MORB Fe8 and Na8 (i.e. FeO and
Na2O values corrected for fractionation to a constant
MgO value of 8 wt %) with ridge axial depth on a global
scale. These correlations have been interpreted, following the experimental data of mantle peridotite melting
(Jaques & Green, 1980), as reflecting varying extent and
pressure of mantle melting in response to mantle solidus
temperature variation of up to 250 K on a global scale.
Langmuir et al. (1992) further elaborated quantitatively
that the global MORB Fe8 variation can be effectively used to calculate the mantle solidus temperature
(T[solidus]) and pressure (P[solidus]) and mantle potential temperature (TP) from the single MORB Fe8 parameter: P[solidus] ¼ 611Fe8 345 (kbar), T[solidus] ¼ 1150 þ
13P[solidus] ( C), and hence TP ¼ T[solidus] P[solidus] 18
( C) by using the experimentally determined dry mantle
solidus curve and by assuming an adiabatic thermal gradient of 18 K kbar–1 (see McKenzie & Bickle, 1988; also
see Niu & O’Hara, 2008).
Because of its convenience, Fe8 has been popularly
and axiomatically used in basalt studies. Given the welldocumented mantle compositional heterogeneity for all
major elements including FeO (e.g. Langmuir &
Hanson, 1980; Niu et al., 1999, 2002a), and to conscientiously inform readers of the inevitable errors when indiscriminately using Fe8 in MORB studies, Niu & O’Hara
(2008) demonstrated in detail in terms of straightforward petrological concepts, physical principles and logical arguments that Fe8 is a misguiding parameter and
cannot be used to infer mantle conditions and processes. In defence of the ‘usefulness’ of Fe8, Langmuir
and his group recently published two papers (Dalton
et al., 2014; Gale et al., 2014) to refute the demonstrations by Niu & O’Hara (2008).
In commemoration of Michael J. O’Hara in this thematic O’Hara volume, and given the fundamental importance of understanding MORB petrogenesis as the
foundation for igneous petrology in particular and global mantle dynamics in general, there is a pressing need
to demonstrate explicitly and objectively that using Fe8
in MORB studies can be misleading and must be treated
with caution. The first-order global MORB major element compositional variation is largely source inherited
and the signal of varying extent of melting as a function
of plate spreading rate variation is also conspicuous as
demonstrated previously (Niu & Hékinian, 1997a).
A more comprehensive and thorough demonstration on
Journal of Petrology, 2016, Vol. 57, No. 11&12
2083
the origin of global MORB chemical systematics will be
presented elsewhere.
THE PROBLEMS OF THE Fe8 PARAMETER
Fe8 is inappropriate in discussing mantle
conditions and processes
It is logical to use basaltic melts in equilibrium with mantle mineralogy (e.g. olivine) to infer mantle melting conditions and processes. This requires that the melt have
Mg# [¼Mg/(Mg þ Fe2þ)] 072 to be in equilibrium with
mantle olivine with Fo content [Mg/(Mg þ Fe)] 090
(or Fo90) because of the well-established olivine–melt
partitioning relationship Mg# ¼ 1/[(1/Fo 1)/Kd þ 1], where
Kd ¼ 030 6 003 at low pressure (Roeder & Emslie, 1970)
or higher values under mantle melting conditions with
primitive MORB compositions (10–15 wt % MgO;
Kd 032; Baker & Stolper, 1994) or primitive ocean island
basalt (OIB) compositions (15–25 wt % MgO; Kd 034;
Matzen et al., 2011).
In Fig. 1 the Klein & Langmuir (1987) data show that
the MORB melts with varying Fe8 values have Mg#
055–068, which is significantly lower than the
Mg# 072 required to be in equilibrium with mantle
olivine. Hence, in terms of the straightforward petrological concept, Fe8 cannot be used to infer mantle processes simply because it records dominantly the highly
evolved signature of MORB melts that have ascended
across the Moho and left the mantle. The mantle solidus
temperatures (T[solidus]) and potential temperatures (TP)
calculated by using Fe8 following Klein & Langmuir
(1987) and Langmuir et al. (1992) thus have no significance. The varying low Mg# values exhibited by MORB
melts with varying Fe8 (as in Fig. 1) are consistent with
varying extents of crustal-level MORB melt evolution
(O’Hara, 1968a, 1968b; Walker et al., 1979; Niu, 1997).
In this context, it is important to note that the popular
emphasis on mantle pyroxenites as sources of OIB and
some alkali basalts (e.g. Sobolev et al., 2007) has led
some to consider that mantle melts generated by pyroxenite melting should have low Mg# (e.g. Yang & Zhou,
2013). This is wrong in concept and incorrect in practice,
for the following reasons: (1) melting of pyroxenites with
Mg# 090 would still produce primitive melts with
Mg# 072 constrained by the Fe–Mg exchange Kd
(i.e. Kdclinopyroxene/melt Kdorthopyroxene/melt Kdolivine/melt 030 6 003 (e.g. Grove et al., 1992; Niu et al., 2002b); (2)
partial melts of pyroxenites with low Mg# values (e.g.
subducted ocean crust) will readily reach chemical equilibrium with the ambient mantle mineralogy to gain
Mg# 072, because pyroxenites are volumetrically
minor; (3) the melts generated by pyroxenite total melting will have the same Mg# as the pyroxenites and will
again reach equilibrium with the ambient mantle mineralogy to gain Mg# 072, regardless of the actual pyroxenite compositions. Hence, ascending mantle melts
prior to crossing the Moho should all be in equilibrium
with mantle mineralogy and have Mg# 072 unless
Fig. 1. (a) Location-averaged global MORB data of Klein &
Langmuir (1987) (KL87) indicating the first-order MORB major
element Fe8 variation (weight per cent FeO corrected for fractionation effect to MgO ¼ 80 wt %) as a function of ocean ridge
axial depth (water depth). Klein & Langmuir used Fe8 to infer
mantle temperature variation (see Langmuir et al., 1992). Niu &
O’Hara (2008) demonstrated that this Fe8 usage is erroneous
because to reveal mantle melting conditions and processes,
MORB melts must have Mg# ¼ Mg/(Mg þ Fe2þ) (10% total Fe is
assumed to be Fe3þ in the Mg# calculation) 072 to be in equilibrium with mantle olivine of Fo 90. However, MORB Fe8 values correspond to Mg# ¼ 054–068 as in (b), which are too
evolved to be in equilibrium with mantle olivine and thus cannot be used to infer mantle melting conditions and processes,
but largely tell us about crustal-level differentiation processes.
(c) combines (a) and (b) to illustrate the variably evolved nature
of MORB melts corrected to Fe8. This figure is simplified from
Niu & O’Hara (2008).
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Journal of Petrology, 2016, Vol. 57, No. 11&12
significant cooling and olivine crystallization has taken
place during melt ascent through the cold thermal
boundary layer beneath ridges (or lithospheric mantle
elsewhere) as seen in abyssal peridotites (Niu, 1997; Niu
et al., 1997). Therefore, if one wants to unmistakably decipher mantle processes from MORB major element
compositions, it is required to correct for the effects of
crustal-level processes to Mg# 072 (e.g. Si72, Ti72,
Al72, Fe72, Mg72, Ca72 and Na72), which must be constrained by the most primitive MORB glass samples
with Mg# 072 (see Niu et al., 1999; Niu & O’Hara,
2008).
The Klein & Langmuir (1987) correction to Fe8
has no significance
Figure 2 is modified from Niu & O’Hara (2008) to compare the fractionation-corrected data of Klein &
Langmuir (1987) with fractionation-uncorrected average
MORB data with MgO 7 wt %. This illustrates that the
apparently meticulous correction procedure of Klein &
Langmuir (1987) is an unnecessary data manipulation
because the corrected Fe8 and Na8 values have no geological difference from the uncorrected data. This is true
for Na2O and largely true also for FeO and Mg# because
the Klein & Langmuir (1987) interpretations were
largely based on the large range (i.e. in Fe8) defined by
samples from the shallowest and deepest ridges (see
Niu, 1997). This evident problem remains despite the
apparently exhaustive arguments in defence of the Fe8
parameter by Gale et al. (2014) as discussed below.
The Gale et al. (2014) justification of Fe8
correction is unjustified
Gale et al. (2014) made a commendable effort in providing an updated MORB dataset for improved studies.
They discussed global MORB systematics and their origin. However, the very heart of the study by Gale et al.
(2014) is to refute the Niu & O’Hara (2008) criticism
against using Fe8 and to defend the Fe8-based interpretations by Klein & Langmuir (1987) and Langmuir et al.
(1992). The Niu & O’Hara (2008) criticism of using Fe8
has three specific points: (1) Fe8, representing highly
and variably evolved MORB compositions, cannot be
used to decipher mantle processes as discussed above;
(2) MgO and FeO are both sensitive and proportional to
the pressure of melting in primitive melts in equilibrium
with mantle mineralogy (see Niu & O’Hara, 2008, fig. 1),
but the use of the Fe8 parameter with constant
MgO ¼ 8 wt % has ignored MgO as a potential pressure
indicator; (3) SiO2 is also sensitive to pressure of melting, but is inversely proportional to MgO and FeO in
primitive melts in equilibrium with mantle mineralogy
(see Niu & O’Hara, 2008, fig. 12a), but MORB SiO2
shows no pressure signature (see Niu & O’Hara, 2008,
fig. 12b and c). Gale et al. (2014) skilfully replied to these
critical points through an apparently painstaking effort,
which is in fact unjustified and has an additional misguiding effect.
Fig. 2. Comparison of location-averaged global MORB Na8 (a),
Fe8 (b) and Mg# (c) [calculated from (b)] with uncorrected
depth-interval averages of Na2O, FeO and Mg# for samples
with MgO > 70 wt % from Niu & O’Hara (2008) (NOH08) to
show that the apparently meticulous correction procedure of
Klein & Langmuir (1987) (KL87) has little significance because
both corrected and uncorrected data are essentially the same,
dominated by a variably evolved nature with Mg# ¼ 054–068.
This figure is simplified from Niu & O’Hara (2008), where details of depth-interval averages are explained in their fig. 3 and
accompanying text.
The Gale et al. (2014) global MORB data in the MgO–
FeO space are illustrated in Fig. 3a. It is evident that the
corrected Fe8 values must be equal to or greater than
the range of 7–125 (55 Fe units) as indicated for FeO
at MgO ¼ 8 wt %. To argue for Fe8 to be useful in reflecting mantle-melting conditions as claimed, Gale et al.
(2014) also corrected MORB melts to be in equilibrium
with mantle olivine of Fo90, denoted as Fe90. As shown
Journal of Petrology, 2016, Vol. 57, No. 11&12
2085
Fig. 3. (a) MgO–FeO covariation plotted using the Gale et al. (2014) MORB dataset, in which the concept of Fe8 is shown by FeO values at MgO ¼ 8 wt % as indicated. (b) The essentially identical Fe8 and Fe90 values and ranges calculated by Gale et al. (2014) were
used to make the case that MORB compositions with Fe8 are the same as MORB Fe90 and thus should be in equilibrium with mantle
olivine of Fo90. However, this is not true because Mg# ¼ f(MgO ¼ 8, Fe8) 054–070, variably and significantly lower than
Mg# ¼ 072 as indicated, showing that Fe8 ¼ Fe90 is conceptually contradictory in terms of elementary petrology. (c) Plot of MgO
against Mg90 and Fe90 from Gale et al. (2014), revealing that the Fe90 values were calculated by adding olivine to each MORB sample until Fe90 ! Fe8, hence obtaining the same Fe8 and Fe90 as in (b); this procedure involves adding as much MgO as required (on
average a factor of two; i.e. Mg90 vs MgO along the horizontal axis), while lowering FeO as needed [lower mean and range of Fe90
than actual FeO in (a)]. (d) Comparison of raw data as in (a) with Gale et al. (2014) calculated Fe8 (the same as Fe90) and Mg90 in
MgO–FeO space, revealing that the Gale et al. (2014) calculated Fe8(Fe90)–Mg90 values (heavily overlapped open red circles) are
remote from the actual data, beyond the petrologically permitted interpolation and extrapolation range. Thus, the Gale et al. (2014)
Fe90–Mg90 values and ranges have no relevance to the actual data.
in Fig. 3b their calculated Fe8 and Fe90 are essentially
the same, with the calculated Fe8 (or Fe90) in the range
of 6–13 (7 Fe units). The significant Fe8–Fe90 correlation with the same range of values looks convincing,
but is curiously incomprehensible. It is puzzling and
confusing why MORB melts are in equilibrium with
mantle olivine Fo90 if expressed with Fe90 (i.e.
horizontal-axis Mg#melt 072–073), whereas these
same MORB melts are not in equilibrium with mantle
olivine if expressed with Fe8 (i.e. vertical-axis
Mg#melt 070 – < 054) despite the simple fact that
Fe8 Fe90 (Fig. 3b). I leave this petrological impossibility
to readers to comprehend, but only offer my personal
understanding of the Gale et al. (2014) correction in
terms of a simple mathematic treatment.
It is obvious that the only way to produce the correlated values in Fig. 3b is to set a given Fe8 value as the
target to produce Fe90 ! Fe8. This can be done by raising MgO while lowering FeO as much or as little as
needed. This is illustrated in Fig. 3c, where MgO must
be raised, on average, by a factor of two (e.g. Mg90)
higher than the actual data (the horizontal-axis values),
whereas FeO must be lowered from the range of 7–16
(horizontal-axis value of Fig. 3a) to Fe90 5–12 (Fig. 3c).
This may be assumed to be equivalent to adding olivine
Fo90 (or other Fo values) until Fe90 Fe8, by adding
more olivine (hence more MgO) to samples with higher
Fe8 and less olivine (hence less MgO) to samples with
lower Fe8. In doing so, the primary melt in the mantle
source region parental to each of the MORB samples
would have Mg# 072–073 in equilibrium with the
mantle mineralogy (see Fig. 3d).
There has been a long history in petrology of adding
olivine to evolved basaltic melts to infer their primary
magma compositions in equilibrium with the mantle
mineralogy. Despite the various uncertainties in doing
so, this is nevertheless a useful approach. The question
is if this approach by Gale et al. (2014) is appropriate
here. The answer is simply ‘No’ because (1) adding
varying amounts of olivine with the purpose of making
Fe90 ! Fe8 is an unjustified exercise, and (2) the so calculated FeO (Fe90, same as Fe8) and MgO (Mg90) in equilibrium with mantle olivine of Fo90 shown by the
stacked red open circles with large MgO (8–19 wt %)
and FeO (6–12 wt %) variations (Fig. 3d) are remotely
removed from, and having no relevance to, the actual
2086
Journal of Petrology, 2016, Vol. 57, No. 11&12
data (the black open circles with MgO 105 wt %).
That is, the Gale et al. (2014) correction to Fe90 is entirely unconstrained and is petrologically unlikely - way
beyond the data and beyond petrologically possible
and permitted interpolation and extrapolation range.
Their ‘olivine addition’ procedure apparently answered the criticisms by Niu & O’Hara (2008) (see
above) because olivine controls the positive Fe90–Mg90
correlation (red circles in Fig. 3d) and negative Si90–
Mg90 and Si90–Fe90 correlations (not shown here) as if
the varying Fe90, Mg90 and Si90 values were consistent
with mantle melting pressure effects. It is objectively
correct to state that MORB melts preserve no melting
pressure signature at all in terms of SiO2 as demonstrated in fig. 12 of Niu & O’Hara (2008) [also see Niu
et al. (2011) and further illustration below]. The correlated Fe90–Mg90–Si90 variations by Gale et al. (2014)
have no pressure significance, but are the artefact of
making Fe90 ! Fe8 (Fig. 3b).
THE EFFICACIES OF MORB MAJOR ELEMENTS
AT Mg# 5 072
About Mg# ‡ 072 and MORB melt compositions
at Mg# 5 072
Basaltic melts having Mg# 072 do not constrain mantle solidus conditions (i.e. Po and To) because Mg# ¼
f(MgO, FeO). A primitive MORB melt with MgO ¼
105 wt % and FeO ¼ 81 wt % gives Mg# 072 (assuming 90% total Fe as Fe2þ) and a primitive ocean island
basaltic melt erupted on thickened lithosphere with
MgO ¼ 18 wt % and FeO ¼ 139 wt % can also give
Mg# 072 (Niu et al., 2011). However, variations of
MORB melt compositions corrected to Mg# ¼ 072 (i.e.
Fe72 and Mg72 as well as Si72, Ti72, Al72 Ca72 and Na72)
can be used to discuss mantle sources and processes
because primitive melts having Mg# 072 are in equilibrium with mantle olivine of Fo 090 (Fo90).
However, MORB melt FeO contents corrected to
MgO ¼ 8 wt % (i.e. Fe8 with Mg# ¼ 056–068) are too
evolved to be in equilibrium with mantle olivine and
cannot be used to discuss mantle processes as elaborated above. The Gale et al. (2014) corrected Fe90 cannot
be used either because Fe90 ¼ Fe8 (Fig. 3b).
Petrological effectiveness of Fe72 versus
misguiding Fe8
Gale et al. (2014) asserted that their new data are not in
agreement with the presentation of Niu & O’Hara
(2008), ‘whose results relied on an inaccurate fractionation correction procedure, which led them to large
errors for high- and low-FeO magmas’ (p. 1051). This is
an incorrect statement as further demonstrated below
in terms of basic petrological concepts and elementary
illustrations.
Figure 4a is a familiar diagram showing global
MORB melt (glass) major element oxide abundance
variations as a function of MgO. Because MgO is
Fig. 4. (a) Global MORB major element oxides [TiO2, Al2O3,
FeOt (where t indicates total Fe expressed as FeO), CaO and
Na2O] abundance variations as a function of MgO (data from
PETDB database: www.petdb.org). (b) As in (a), but plotted as a
function of Mg# [Mg/(Mg þ Fe2þ), assuming Fe2þ/total Fe ¼ 09
for simplicity]. As MgO is linearly proportional to the liquidus
temperature (e.g. Weaver & Langmuir, 1990; Danyushevsky,
1998; Niu et al., 2002b; Niu & O’Hara, 2009), all the oxide variations with decreasing MgO define first-order liquid lines of
descent (LLD) during cooling-dominated MORB melt cooling
and evolution. In detail, however, these trends are not simple
LLD, but the combined effect of LLD, magma mixing, melt–rock
reactions and complex open-magma chamber processes
(O’Hara, 1977). In other words, these trends are the net effect
of all these natural processes. Mg# is not strictly linearly proportional to the liquidus temperature, but is convincingly close
for MORB melts with MgO > 4 wt % (see Niu et al., 2002b;
Stone & Niu, 2009), with deviation caused by varying FeO content. Hence, to a first order, MORB major element oxide abundance variations versus Mg# in (b) are also a net effect of the
combined MORB melt evolution trends. As indicated by the
vertical gray dashed lines, the band width for each oxide would
be the values and ranges at MgO ¼ 80 wt % in (a) or at
Mg# ¼ 072 in (b), respectively.
proportional to the liquidus temperature, the data trend
for each of the oxides with decreasing MgO is to a first
order consistent with the liquid line of descent (LLD),
which is conceptually the residual melt compositional
variation as the result of fractional crystallization of
olivine þ spinel ! plagioclase þ olivine ! plagioclase þ
Journal of Petrology, 2016, Vol. 57, No. 11&12
clinopyroxene 6 olivine prior to reaching MgO 40 wt %,
after which crystallization of titanomagnetite (Ti–Fe
oxides) depletes FeO and TiO2 while increasing SiO2,
entering the basaltic-andesite stage of basaltic magma
evolution (e.g. Niu et al., 2002b; Niu, 2005a; Niu &
O’Hara, 2009; Stone & Niu, 2009). Strictly speaking, the
data trends in Fig. 4a are not simple LLD, but the net effect of cooling-dominated crustal-level processes combined (e.g. fractional crystallization, magma mixing,
melt–rock assimilation and, or, reaction and other aspects of complex open-magma chamber processes).
The Klein & Langmuir (1987) and Langmuir et al.
(1992) method of calculating Fe8 and Na8 was to assume a common LLD slope for FeO–MgO and Na2O–
MgO, respectively, to project individual data points
onto the plane of MgO ¼ 8 wt % as indicated, by backtracking (for samples with MgO < 8 wt %) and forwardtracking (for samples with MgO > 8 wt %). To avoid
correction errors, Klein & Langmuir (1987) and
Langmuir et al. (1992) chose samples with
MgO 55 wt %. One can do the same to obtain Ti8,
Al8, Ca8 and Ca8/Al8 (Niu & Batiza, 1991a, 1993, 1994).
The only difference of the Gale et al. (2014) approach
from that of Klein & Langmuir (1987) and Langmuir
et al. (1992) was to obtain separate sets of LLD for
MORB samples from each ridge segment with the belief that samples from different ridge segments have
different LLD trends as discussed by Niu & Batiza
(1993). This philosophical consideration is reasonable,
but has unavoidable problems: (1) MORB samples
from most ridge segments do not have enough MgO
variation coverage to obtain suitable LLD (if they were
LLD); hence, assumptions must be made to justify the
chosen LLD slopes that in practice cannot be constrained; (2) natural MORB sample suites do not define
pure LLD, but show the combined effects of complex
crustal-level processes (see above).
The same data as a function of Mg# are plotted in
Fig. 4b. The data trends are again not simple LLD, but
the net effect of compound crustal-level processes (see
above). To correct for such net effect for the global dataset, it is logical to find a common set of correction coefficients for each oxide, applicable to the global dataset
in this Mg# variation diagram. Obviously, the most objective, logical and simplest method is to obtain a set of
polynomial regression coefficients for each oxide based
on the entire global dataset (see Niu et al., 1999; Niu &
O’Hara, 2008) and to project each sample along the
polynomial curves to Mg# ¼ 072 as indicated by the
vertical gray dashed line in Fig. 4b. To avoid unnecessary errors, Niu & O’Hara (2008) used MORB samples
with MgO 70 wt %, thus the polynomial coefficients
mostly give rise to linear coefficients, especially for FeO
and MgO [see Appendix of Niu & O’Hara (2008)]. This
method of correction is objective, logical and simple for
the following reasons.
1. We follow the data (Fig. 4b) without making any unnecessary assumption that cannot be justified.
2087
2. Regardless of how evolved and how complex the
histories each MORB sample may have experienced
through crustal-level processes, its parental melt
must have once had Mg# 072 within the range
constrained by the data as indicated by the gray
dashed line at Mg# ¼ 072 (Fig. 4b).
3. The method corrects for the net effect of compound
crustal-level processes expressed by the global data
trends without having to know exactly what LLD for
a single sample suite may be (Fig. 4b).
4. Mg# ¼ 072 is the best reference point: (a) the most
primitive MORB samples have Mg# 072–073 (corresponding to some of the samples with
MgO 105 wt % in Fig. 4a) so that the correction is
well constrained within the known dataset (Fig. 4b)
without having to make assumptions that cannot be
justified; (b) Mg# ¼ 072 is about the minimum value
for a basaltic melt required to be in equilibrium with
mantle olivine.
5. The key difference of the Niu & O’Hara (2008) correction method is not an ‘inaccurate fractionation procedure’ as incorrectly stated by Gale et al. (2014), but
the use of the petrologically and logically best reference of Mg# ¼ 072 (Fig. 5b) rather than the problematic reference of MgO ¼ 80 wt % (Fig. 5a).
We should note that correction to MgO ¼ 8 wt % or
Mg# ¼ 072 has less effect on the relative abundances
and systematics of Ti72 (vs Ti8), Al72 (vs Al8), Ca72 (vs
Ca8) and Na72 (vs Na8), as well as Ca72/Al72 (vs Ca8/Al8),
but has large effects on Fe72 (vs Fe8) and Mg72 (vs
Mg8 ¼ 8) because of the FeO–MgO–Mg# interrelationships as further illustrated below.
More on the effectiveness of Fe72 versus
misguiding Fe8 (and Fe90)
The variation of FeOt against MgO and Mg# respectively (Fig. 5a and b) explicitly shows that Fe72 is objective and more logical than Fe8, and why Fe72 values
must have a significantly smaller range (10 Fe unit)
than Fe8 (45 Fe units) without any correction. With
decreasing MgO (or Mg#), the upper bound of FeOt
reaches the maximum at MgO 4 wt % (or Mg# 030)
as a result of combined fractional crystallization of spinel þ olivine þ plagioclase þ clinopyroxene before Fe–Ti
oxide (titanomagnetite) appears on the liquidus at
MgO 4 wt % to cause FeOt depletion. By considering
samples with MgO > 55 wt % only, this can certainly
avoid potential errors caused by highly evolved samples in correcting to Fe8 (Klein & Langmuir, 1987;
Langmuir et al., 1992; Gale et al., 2014). However, the
actual data, with no correction, show that at MgO ¼ 8 wt
%, FeO ¼ Fe8 72 117 wt % has an Fe8 range of
45 Fe units as indicated (Fig. 5a). This is the very minimum Fe8 spread because correction for samples with
MgO > 55 wt % and < 80 wt % onto the MgO ¼ 80 wt %
plane can only make the Fe8 spread larger (i.e. > 45
Fe units). On the other hand, it is evident that FeO at
or projected to Mg# ¼ 072 has a limited variation of
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Journal of Petrology, 2016, Vol. 57, No. 11&12
Fig. 5. The same two panels as in Fig. 4, but only FeOt is plotted to highlight the difference between Fe8 (a) and Fe72 (b). The maximum of
the FeOt upper bound at MgO 4 wt % in (a) is to a first order consistent with a liquid line of descent resulting from the combined effects of
fractional crystallization of spinel þ olivine þ plagioclase þ clinopyroxene before Fe–Ti oxide (titanomagnetite) on the liquidus at
MgO 4 wt %. Klein & Langmuir (1987) (KL87), Langmuir et al. (1992) (LKP92) and Gale et al. (2014) (GLD14) used samples with
MgO 55 wt % and Niu & O’Hara (2008) (NOH2008) used samples with MgO 70 wt %, so the obviously scattered FeO values below these
chosen cut-off MgO values will not affect the calculated Fe8 or Fe72. However, those samples below the main trend in the broad shaded region cannot simply be evolved MORB melts, but are most consistent with being ‘snapshots’ of magma mixing in open magma chamber (or
something equivalent) systems, largely resulting from mixing of newly replenished primitive melts with already highly evolved melts having
low MgO and low FeO. There is no doubt that many samples towards the lower bound of the main data band at a given MgO or Mg# have
resulted from such mixing, although it is not possible to precisely quantify this process. It is reasonable to infer that such mixing is more obvious in FeO–MgO space (a) than in the FeO–Mg# space (b) because of the significantly fatter lower bound at a given MgO in (a). The key message here is as follows: for this same dataset, Fe8 7–115 wt % with variation of 45 Fe units, corresponding to Mg# ¼ 068 – 057, far too
evolved to be used for discussing mantle melting conditions and processes. On the other hand, Fe72 7–80 wt % with 10 Fe unit variation
at a constant Mg# ¼ 072, which is in equilibrium with mantle olivine of Fo90 and thus can be logically used to discuss mantle melting conditions and processes. It is straightforward that the small Fe72 range of 1 Fe unit [versus the large Fe8 range of >45 Fe units found by Gale
et al. (2014)] is a property of the actual data, and is not the result of an inaccurate correction procedure as incorrectly stated by Gale et al.
(2014).
Journal of Petrology, 2016, Vol. 57, No. 11&12
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7–8 wt % with an Fe72 range of 1 Fe unit as illustrated. Given the narrow FeO band at a given Mg#, and
because all the samples, regardless of how they
evolved, were modified through complex crustal-level
processes, their parental melts must have once passed
through Mg# ¼ 072 with an Fe72 spread that cannot be
significantly greater than that observed. This is a
straightforward demonstration based on the actual data
without applying any correction, in contradiction to the
incorrect statement by Gale et al. (2014).
In summary, the data, not the correction procedure,
demonstrate that Fe8 72 117 wt % with a large range
of 45 Fe units (Fig. 5a; Klein & Langmuir, 1987;
Langmuir et al., 1992; Gale et al., 2014). Also the data, not
correction
procedure,
demonstrate
that
Fe72 ¼
70 80 wt % with a small range of 1 Fe unit (Fig. 5b;
Niu & O’Hara, 2008). This ‘coincidence’ of the actual FeO
at Mg# ¼ 072 with Fe72 by Niu & O’Hara (2008) demonstrates the objective validity of the Niu & O’Hara (2008)
correction procedure, which strictly follows basic petrological principles and the intrinsic structure within the data.
Figure 6a objectively compares Fe72 (Niu & O’Hara, 2008)
with Fe8 (Klein & Langmuir, 1987; Langmuir et al., 1992) in
Na2O–FeO space. Figure 6b and Fig. 6c are the same, but
using the Gale et al. (2014) data to compare the Niu &
O’Hara (2008) method (250 m ridge depth interval averages) with the 237 global ridge segment averages of Gale
et al. (2014) (excluding the three averages from Iceland on
land). It should be noted that Niu & O’Hara (2008) used
9130 MORB glass samples with MgO > 70 wt %, but there
are only 7607 samples with MgO > 70 wt % and no data
for the depth interval of 4750–5000 m in the Gale et al.
(2014) dataset. Nevertheless, Fig. 6b (comparison of Na72
and Fe72 with Na8 and Fe8) and Fig. 6c (comparison of
Na72 and Fe72 with Na90 and Fe90) are essentially the same
as Fig. 6a in terms of process interpretations. Therefore,
the statement by Gale et al. (2014) against the Niu &
O’Hara (2008) method is inappropriate, and has the effect
of further misguidance in addition to the misleading Fe8.
The proposed large mantle potential temperature variation
of 250 K on the basis of the large Fe8 range [and the
same large Fe90 range obtained by the Gale et al. (2014)
method] is manifestly an artefact. Therefore, any intended
use of Fe8 should be avoided in order to advance our
genuine understanding of ocean ridge dynamics.
THE MEANING OF MORB MAJOR ELEMENTS AT
Mg# 5 072
We can now use MORB major element compositions at
Mg# ¼ 072 to infer mantle processes. Because ridge
axial depth variation (0 to 6000 m) and plate spreading rate variation (<10 to >150 mm a–1) are the only two
largest known physical variables along the global ocean
ridge system, our task is to examine whether MORB
major element compositions respond to these two
physical variables. Klein & Langmuir (1987) recognized
Na8 and Fe8 correlations with ridge axial depth and Niu
& O’Hara (2008) confirmed this but emphasized that
Fig. 6. (a) Comparison of Na8–Fe8 of Klein & Langmuir (1987)
(KL87) with Na72–Fe72 of Niu & O’Hara (2008) (NOH08). Na8 and
Na72 values and ranges are similar, which is obvious from Fig.
4, but Fe8 and Fe72 are very different, with variations in Fe8 5
units compared with Fe72 1 unit, which is rather obvious in
Figs 4 and 5. This straightforward difference between Fe8 and
Fe72 as demonstrated by the actual data (Figs 4 and 5) is an intrinsic property of the data, but was incorrectly interpreted by
Gale et al. (2014) as an effect of an erroneous or inaccurate
fractionation correction procedure of Niu & O’Hara (2008). (b,
c) As in (a), the same comparison of Na72–Fe72 using the Gale
et al. (2014) (GLD14) dataset with the Na8–Fe8 of Gale et al.
(2014) (b) and Na90–Fe90 from Gale et al. (2014) (c). Samples
from Iceland and nearby ridges with seawater depths < 200 m
are excluded.
2090
Ti72, Al72, Mg72, Ca72, Ca72/Al72 as well as Na72 and Fe72
all correlate with ridge axial depth and offered different
explanations for the cause of such correlations. Niu &
Hékinian (1997a) demonstrated the significant correlations of Al8 and Ca8/Al72 (also K8/Ti8 as shown below)
with plate spreading rate, which, together with the then
available global abyssal peridotite data, were interpreted to reflect the extent of mantle melting, which increases with increasing plate spreading rate as expected
in terms of straightforward physics (Reid & Jackson,
1981; Niu & Hékinian, 1997a). However, Gale et al.
(2014) denied the existence of such a correlation by stating: ‘There is no correlation between the chemical parameters and spreading rate’ (p. 1080). In this section, I
discuss these correlations, but it is conceptually important first to clarify that spreading rate variation and ridge
depth variation are two genetically unrelated variables
(except for extremely slow-spreading ridges; see below)
and we need to separate the effects of the two variables
for a better understanding of ridge processes.
The origin of plate spreading rate variation and
ridge axial depth variation
Plate spreading rate here refers to the separation rate of
the two plates with respect to each other on both sides of
the ridge, which is also called the ‘full spreading rate’ or
‘total opening rate’ of the ridge, and is determined by the
absolute motion speeds of the two plates relative to some
hotspot reference framework (e.g. DeMets et al., 1990).
The absolute motion speed of a given plate has been well
understood as resulting from subducting slab pull (e.g.
Forsyth & Uyeda, 1975; McKenzie & Bickle, 1988; Davies
& Richards, 1992), which is straightforward to understand
for plates connected with subduction zones such as those
in the Pacific, but not so obvious for plate motion in the
Atlantic. Nevertheless, Niu (2014, 2016) illustrated that (1)
seafloor spreading in ocean basins with passive margins
(e.g. the Atlantic type) and (2) continental drift are simply
passive movements in response to trench retreat of active
seafloor subduction in ocean basins like the Pacific with
subduction zones. Hence, the varying speeds of plate motion determine the plate spreading rates of the ridges concerned. Because ocean ridge magmatism results from
plate separation, it is expected that the effect of plate separation rate variation must be recorded in the MORB
chemistry. To deconvolve the effect of plate spreading
rate variation from the effects of all other probable or possible variables involves averaging MORB chemistry with
respect to spreading rate intervals, so as to average out all
other possible effects (Niu & Hékinian, 1997a). This is
mathematically the same as seeking the partial derivatives
of a quantity with respect to a particular variable, with all
other variables held constant (Niu & O’Hara, 2008).
The above elaboration demonstrates that globally
the large ridge axial depth variation is unrelated to
spreading rate (see below for very slow spreading
ridges). This is because ridge axial depth variation is
largely controlled isostatically by sub-ridge material
Journal of Petrology, 2016, Vol. 57, No. 11&12
density variations owing to varying thermal or compositional buoyancy variations that are unknown to slab pulls
in remote subduction zones. For example, the pulling of
the subducting slabs in the western Pacific or beneath the
Andes does not know the sub-ridge buoyancy and ridge
axial depth variation at the southern East Pacific Rise.
Nevertheless, it has been known for decades that ridge
spreading rate differences do control across-ridge morphology and along-ridge topography on local and ridge segment scales, especially at slow-spreading ridges (e.g.
Macdonald, 1982; Macdonald et al., 1988; Sempere et al.,
1990; Sinton et al., 1991; Grindlay et al., 1991; Niu & Batiza,
1993; Niu et al., 2001; Dick et al., 2003). These local and
segment-scale variations of ridge depth and morphology
are most probably of lithospheric origin and isostatically
uncompensated (i.e. dynamic topography or morphology).
The effects of these features, if any, thus must be removed
to reveal the origin of the correlated variations of MORB
chemistry with ridge axial depth on a global scale. Klein &
Langmuir (1987) and Gale et al. (2014) used regionally
smoothed ridge axial depth to obtain an average depth
value for a given ridge segment. This requires assumptions on how to smooth and how to choose a segment
depth value. Instead, Niu & O’Hara (2008) made no assumptions, but used actual MORB sample depths [see
below and also fig. 3 of Niu & O’Hara (2008)].
The meaning of MORB chemistry–ridge depth
correlation
As above for elucidating the spreading rate effect, to effectively deconvolve the effect of processes that lead to
ridge axial depth variations from the effects of all other
probable or possible variables, it is desirable to average
MORB chemistry with respect to ridge axial depth intervals. The averaging is not arbitrary, but follows basic
principles (Niu & O’Hara, 2008), as listed here.
1. A 250 m depth interval is chosen because this depth
interval size approximates the ridge depth variation
over 500 km regional scale-lengths. This gives >20
average data points with which to work. This number
of data points is large enough to be statistically significant without compromising first-order systematics, but
small enough to smooth out secondary effects while
making the first-order global systematics prominent.
2. The averaging uses actual sample depths regardless
of geographical location and ocean basin (i.e. the
Pacific, Atlantic and Indian Oceans).
3. Such heavy averaging averages out the effects of (a)
spreading rate variation, (b) local-scale fertile source
compositional heterogeneity [including the arbitrary
normal (N-) and enriched (E-)type MORB division],
and (c) dynamic topography on regional and ridge
segment scales (see above) to objectively observe
the significance of MORB chemistry variation with
respect to ridge axial depth.
Figure 7 is modified from Niu & O’Hara (2008). The
22 data points represent heavy averages of 9130 global
Journal of Petrology, 2016, Vol. 57, No. 11&12
MORB data with MgO 7 wt %, corrected to Mg# ¼ 072
within each of the 22 depth intervals (bins) of 250 m
from actual ridge depths of 250 to 5750 m below sea
level (see Niu & O’Hara, 2008, fig. 3). Figure 7a and b
shows significant correlations of MORB Ti72, Al72, Fe72,
Mg72, Ca72, Na72 and Ca72/Al72 with ridge axial depth.
Such correlations suggest a genetic relationship between MORB chemistry and ridge axial depth.
However, this relationship, cannot be a ‘cause and effect’ one because ridge axial depth (water depth) cannot
physically control MORB chemistry. I consider that the
correlated variations of MORB chemistry with ridge
axial depth in Fig. 7a and b are two different effects of a
common cause. What is this common cause?
MORB chemistry–ridge depth correlation results
from mantle compositional variation
Niu & O’Hara (2008) demonstrated explicitly that the intrinsic cause is the fertile mantle major element compositional variation. This is because fertile mantle major
element compositions determine (1) variation in both
composition and mode of mantle mineralogy, (2) variation of mantle density, (3) variation of ridge axial depth
because of the mantle density variation, (4) sourceinherited MORB compositional variation, (5) densitycontrolled variation in the maximum extent of mantle
upwelling, (6) apparent variation in the extent of melting, and (7) the correlated variation of MORB chemistry
with ridge axial depth. These physically straightforward, logically sound and geologically simple demonstrations are summarized in Fig. 7c.
The physical significance of the correlation between
MORB chemistry and ridge axial depth is shown in Fig. 7
in the form of schematic illustrations, which are convenient for conceptual clarity [see Niu & O’Hara (2008) for
quantitative analysis; their figs 13–19 and related discussion]. It has long been recognized that source rock compositions exert a key control on the resulting magmas,
as summarized by Cox (1992): ‘the most fundamental
thing to understand is that the spectrum of compositions
we see is controlled in the first place by the composition
of the Earth itself. Just as in the simplest phase diagram
the course of evolution of residual liquids by fractionation, or the course of liquid evolution during partial
melting, is controlled by the starting composition, so in
the Earth the range of igneous rocks is constrained by
the original materials’. This simple concept of source
control has been repeatedly demonstrated experimentally (see Jaques & Green, 1980; Green, 2015; Green &
Falloon, 2015), and also well documented observationally (e.g. Mahoney et al., 1994; Niu et al., 1996, 1999,
2001, 2002a; Castillo et al., 1998, 2000). Furthermore, on
the basis of their detailed petrology and geochemistry
study of several ridge segments along the northern MidAtlantic Ridge, Niu et al. (2001) recognized the dynamic
role of mantle source compositional variation: ‘Mantle
compositional control on the extent of mantle melting,
2091
crust production, gravity anomaly, ridge morphology,
and ridge segmentation’.
Indeed, as demonstrated by Niu & O’Hara (2008), the
simplest interpretation is that the MORB major element
compositional systematics illustrated in Fig. 7a and b
largely derive from the fertile mantle compositional systematics. That is, with increasing ridge axial depth,
MORB Ti72, Al72 and Na72 increase whereas Fe72, Mg72,
Ca72 and Ca72/Al72 decrease, reflecting variations in the
fertile mantle source composition. In other words, corresponding to the increasing ridge axial depth, the subridge mantle source increases in TiO2, Al2O3 and Na2O,
and decreases in FeO, MgO, CaO and CaO/Al2O3,
becoming progressively less depleted (or more enriched) in a basaltic melt component. This correspondence becomes physically straightforward if we consider
the source mineralogy as a function of the major element compositions (Niu, 1997). Correlated with increasing ridge depth, the mantle source region is
progressively (1) enriched in pyroxene over olivine (or
higher pyroxene/olivine ratio) because of decreasing
FeO and MgO, (2) enriched in jadeite over diopside (or
higher jadeite/diopside ratio) in clinopyroxene because
of increasing Na2O and Al2O3 and decreasing CaO and
MgO, and (3) enriched in garnet because of increasing
Al2O3. These mineralogical systematics consistently indicate increasing asthenospheric mantle density from
beneath shallow ridges to beneath deep ridges (see Niu
& Batiza, 1991b, 1991c; Niu et al., 2003), which explains
in terms of isostasy why the progressively deeper ridge
is underlain by progressively more enriched (or less
depleted) denser mantle and why there is such MORB
chemistry–ridge depth correlation as demonstrated
quantitatively by Niu & O’Hara (2008) (see their figs 13–
19 and detailed discussion) with the compensation
depth probably in the mantle Transition Zone (Niu &
O’Hara, 2008).
This fertile mantle compositional variation controlled
mantle density variation can in turn have a physical
feedback on ridge dynamics. Because the sub-ridge
mantle upwelling is a passive response to plate separation (McKenzie & Bickle, 1988), and because the extent
of the upwelling (i.e. PoPf in Fig. 7c) determines the extent of decompression melting (F / PoPf), the dense
fertile mantle beneath deep ridges thus has restricted
upwelling (isostatic control), which gives way to conductive cooling to a deep level, forces melting to stop at
such a deep level, leads to a short melting interval, and
thus produces less melt and probably a thin magmatic
crust relative to the less dense (more refractory) fertile
mantle beneath shallow ridges as illustrated in Fig. 7c
(see Niu & O’Hara, 2008). This increased final depth of
melting (Pf), the increased thickness of the ‘cold thermal
boundary layer’ (CTBL), and hence the reduced extent
of decompression melting (F / PoPf) in response to
the increased fertility and density of the MORB mantle
source beneath deep ridges collectively result from a
direct physical control; that is, the well-understood ‘lid
effect’ (Niu et al., 2011). The effect of this physical
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Journal of Petrology, 2016, Vol. 57, No. 11&12
Fig. 7. (a, b) Modified from Niu & O’Hara (2008) to emphasize that, except for SiO2, all the major element oxides (TiO2, Al2O3, FeO,
MgO, CaO, Na2O and CaO/Al2O3), not just FeO and Na2O, of the global MORB data after correction for the effects of cooling-dominated crustal-level processes to Mg# ¼ 072, show significant intercorrelations and systematic correlations with ocean ridge axial
depth. Such correlations are most consistent with fertile MORB mantle source compositional variation. The fertile mantle source is
progressively more enriched (or less depleted) from beneath shallow ridges than beneath deep ridges, pointing to increasing
modal garnet, jadeite/diopside ratio in clinopyroxene, and pyroxenes/olivine ratio, and thus a progressively denser mineral assemblage towards deep ridges as shown in (c). This explains straightforwardly why deep ridges are deep—fertile asthenospheric mantle is denser than more depleted mantle (see Niu et al., 2003), thus resulting in deeper ridges. Shallower ridges are underlain by
more depleted asthenospheric mantle and controlled by isostasy (see Niu & O’Hara, 2008, figs 13 and 14 and accompanying discussion). In addition, dense fertile mantle beneath deep ridges upwells reluctantly in response to plate separation, which leads to a limited extent of upwelling, allowing conductive cooling to penetrate to a great depth, making a thickened cold thermal boundary
layer (CTBL; the gray triangular region labeled ‘2’ in the illustration), forcing melting to stop at a deep level (Pf), resulting in a short
melting interval (PoPf) and less melting, and probably producing a thin magmatic crust relative to the more refractory fertile mantle beneath shallow ridges. The shaded triangular areas in the P–T diagrams illustrate conceptually the difference in melt production that is proportional to the height of the decompression melting intervals (PoPf). Thus, deep ridge MORB have signatures of
apparently lower extents of melting than shallow ridge MORB [see (a) and (b)] superimposed on the fertile mantle compositional inheritance [see details given by Niu & O’Hara (2008)]. An important point here is that globally the solidus depth (Po) beneath ocean
ridges away from mantle plumes (or anomalous mantle with abundant volatiles and alkalis) is essentially constant or varies little
(e.g. To < 50 K if any; see Niu & O’Hara, 2008). In contrast, the final depth of melting (Pf) can vary because of varying thickness of
the CTBL, termed the ‘lid effect’, which is conspicuous for intra-plate ocean island basalts (see Niu et al., 2011), but is also detectable beneath ocean ridges as the result of plate spreading rate variation as demonstrated by Niu & Hékinian (1997a) and further
illustrated in Fig. 8 using the new global dataset of Gale et al. (2014).
Journal of Petrology, 2016, Vol. 57, No. 11&12
feedback on MORB chemistry is expected to be small,
but produces the same systematics as those inherited
from fertile source compositions and thus amplifies the
MORB chemistry correlation with ridge axial depth as in
Fig. 7b and c.
Hence, the global MORB compositional systematics
(Fig. 7b and c) are the net effect of (1) fertile mantle
source inheritance and (2) varying extents of melting
controlled by the varying extent of upwelling and decompression melting as the result of mantle density
variations ultimately still controlled by fertile mantle
compositional variation. Dick and co-workers have recently offered supporting case studies on the effect of
fertile mantle compositional control on MORB chemistry and ridge depth (Zhou & Dick, 2013; Dick & Zhou,
2015).
The overstated mantle temperature variation is
secondary or has negligible effect on the MORB
chemistry–ridge depth correlation
Klein & Langmuir (1987), Langmuir et al. (1992) and
Gale et al. (2014) used the large Fe8 variation (on average 5 Fe units; Figs 1–6) to argue for large mantle temperature variations of 250 K beneath global ocean
ridges (this refers to mantle solidus depth variation,
DPo, and corresponding solidus temperature variation,
DTo) as the cause of the MORB chemistry–ridge depth
correlation. Because Fe8 (and their Fe90 ¼ Fe8; Fig. 3b) is
not a mantle signature, but a highly and variably
evolved crustal-level signature with Mg# ¼ 056–068
that cannot be used to infer mantle conditions in the
first place, there is no petrological evidence for large
mantle potential temperature variations along the global ocean ridges. Given the small thermal expansion coefficient of 3 10–5 K–1 under upper mantle conditions,
the effect of mantle temperature variation on ridge axial
depth variation, if any, is negligible as vigorously demonstrated by Niu & O’Hara (2008) (e.g. 1% density reduction owing to compositional depletion is equivalent
to temperature increase of 330 K). Hence, mantle temperature variation as the cause of ridge depth variation
can be ruled out on the basis of simple mineral physics.
To defend the interpretation of mantle temperature control, Langmuir & co-workers used ridge axial depth as a
constraint for mantle temperature variation (Dalton
et al., 2014); their supporting argument is the low seismic velocity in the deep mantle (>300 km) beneath
Iceland and adjacent shallow ridges. Evidently, using
ridge depth as a mantle temperature constraint is subjective and circular, rather than justified evidence, because this argument neglects the intrinsic control of
fertile mantle compositional variation on the physical
properties and dynamic role of the mantle, as well as on
MORB compositional variation.
In this context, it is conceptually important not to
confuse plate tectonics with mantle plumes because
they have no genetic connection and represent different
modes of Earth’s cooling. It is well understood that
2093
plate tectonics is driven by the top cold thermal boundary layer (lithospheric plates) that cools the mantle,
drives major aspects of mantle convection and explains
seafloor subduction and ridge dynamics (McKenzie &
Bickle, 1988; Davies & Richards, 1992), whereas mantle
plumes are driven by the basal hot thermal boundary
layer (core–mantle boundary) (e.g. Richards et al., 1989;
Campbell & Griffiths, 1990; Davies & Richards, 1992)
despite some debate (e.g. Davies, 2005; Foulger, 2005;
Niu, 2005b). However, when mantle plumes rise to
reach the lithospheric plates, interaction between the
two can take place. Because the lithosphere is the thinnest at ridges, such interaction is best expressed as
plume–ridge interactions (e.g. Ito et al., 2003; Niu &
Hékinian, 2004). To study ridge processes of plate tectonic origin, we should study ridges uninfluenced by
hotspots or mantle plumes. Hence, to use the seismic
low Vs anomaly at >300 km depth beneath Iceland
(Dalton et al., 2014; Gale et al., 2014) that is absent elsewhere beneath global ocean ridges as evidence of a hot
ridge of plate spreading origin and as evidence of large
mantle temperature variations beneath global ocean
ridges is inappropriate. In fact, the invoked large mantle
temperature variation disappears when data from
plume-influenced ridges are removed (see Niu &
Hékinian, 1997a; Niu & O’Hara, 2008; Regelous et al.,
2016; also see below). There is no evidence of any kind
for large mantle temperature variation beneath ridges
away from mantle plumes.
I will specifically discuss in a later section why
MORB melts do not preserve the solidus conditions (i.e.
Po and To) as previously elaborated (Niu, 1997).
The meaning of MORB chemistry–spreading rate
correlation
Because ocean ridges are of plate tectonic origin, and
because sub-ridge mantle upwelling is a passive response to plate separation, the rate of mantle upwelling
must be positively related to plate spreading rate (e.g.
Reid & Jackson, 1981; Phipps Morgan, 1987; Niu, 1997).
This means that the sub-ridge mantle thermal structure
and extent of decompression melting must also vary as
a function of plate spreading rate. Indeed, Niu &
Hékinian (1997a) used the then available global MORB
data (Niu & Batiza, 1993) together with the available
abyssal peridotite data (Dick & Fisher, 1984; Dick et al.,
1984; Dick, 1989; Johnson et al., 1990; Niu & Hékinian,
1997b) to demonstrate that the extent of sub-ridge mantle melting increases with increasing spreading rate as
anticipated, and as recently confirmed with an updated
dataset (Regelous et al., 2016). Figure 8a–c is reproduced from the spreading rate interval (bin width
20 mm a–1) averaged MORB data of Niu & Hékinian
(1997a), plotted as a function of full spreading rate. As
above, heavy averaging is required to remove the effects of ridge axial depth variation and location-related
mantle source heterogeneity variation on varying
scales (including the arbitrary N- and E-type MORB
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Journal of Petrology, 2016, Vol. 57, No. 11&12
Fig. 8. Plots (a)–(c) from Niu & Hékinian [1997a; (c) was not shown then] together with the then available abyssal peridotite data
allowed them to conclude that the extent of mantle melting beneath ocean ridges increases with increasing plate spreading rate.
However, Gale et al. (2014) concluded (p. 1080): ‘There is no correlation between the chemical parameters and spreading rate.’
Plots in (d)–(f) use the new dataset of Gale et al. (2014) (GLD14) to show essentially the same as shown by Niu & Hékinian (1997a);
that is, that decreasing Al72 (d) and Al8 (a), increasing Ca72/Al72 (e) and Ca8/Al8 (b), and decreasing K2O/TiO2 (f) and K8/Ti8 (c) with
increasing spreading rate are simply and straightforwardly the result of increasing extent of melting. This spreading rate dependent
extent of melting is a consequence of the spreading rate controlled lid effect as illustrated in (g), where notations are the same as in
Fig. 7c. The numerals next to the data points in (a) and (d) denote ‘20 mm a–1’ spreading rate intervals for averaging. Details for (a)
and (b) have been given by Niu & Hékinian (1997a). Details in (d)–(f) (the insets show all the data and averages) are as follows:
1, <20 mm a–1, n ¼ 891 samples; 2, 20–40 mm a–1, n ¼ 2163; 3, 40–60 mm a–1, n ¼ 1923; 4, 60–80 mm a–1, n ¼ 1228; 5, 80–100 mm a–1,
n ¼ 1406; 6, 100–120 mm a–1, n ¼ 1130; 7, 120–140 mm a–1, n ¼ 7; 8, 120–140 mm a–1, n ¼ 236. The heavy averaging is necessary to
examine objectively the presence and significance of any spreading rate effect by averaging out all other non-spreading rate factors
(e.g. effects of fertile source variation, including the arbitrary N- and E-type MORB, ridge axial depth variation and uncertainties
associated with the correction procedures). Samples with seawater depth <1500 m are excluded, as are those from plumes and
plume-influenced ridges (these are samples from Iceland, near-Iceland ridges and the Red Sea) as indicated in Fig. 9a. The reason
for choosing <1500 m depth is because most normal ridges have axial depths >2000 m, but there are ridges far away from plumes
or hotspots that can be as shallow as 1500 m; for example, the ridge OH-1 south of the Oceanographer Fracture Zone at 35 N
MAR (Niu et al., 2001).
Journal of Petrology, 2016, Vol. 57, No. 11&12
division) to reveal if there indeed exists any genuine
correlation between MORB chemistry and plate spreading rate. As Al is an incompatible element and K is more
incompatible than Ti during sub-ridge mantle melting,
the decreasing Al8 and K8/Ti8 and increasing Ca8/Al8
with increasing spreading rate are a straightforward expression of an extent of mantle melting that increases
with increasing spreading rate. Figure 8g provides a
straightforward illustration in terms of the lid effect (i.e.
the CTBL thickness) as to why the extent of mantle melting increases with increasing spreading rate. Because
the rate of mantle upwelling increases with increasing
plate spreading rate (see above), fast upwelling beneath
fast-spreading ridges allows the adiabat to extend to a
shallower level against conductive cooling to the surface. By contrast, with slower upwelling beneath slowspreading ridges, conductive cooling to the surface extends to a greater depth against the adiabat. Hence, the
depth of Pf decreases and the decompression interval
(PoPf) increases with increasing spreading rate, hence
the extent of melting F (/ PoPf) increases with increasing plate spreading rate.
However, Gale et al. (2014) refuted the demonstration and interpretation by Niu & Hékinian (1997a) by
stating: ‘There is no correlation between the chemical
parameters and spreading rate’ (p. 1080). Figure 8d–f is
based on the Gale et al. (2014) data and argues against
their unfounded statement, confirming the findings and
demonstration by Niu & Hékinian (1997a) that the extent of sub-ridge mantle melting increases with increasing spreading rate. It is important to compare Fig. 8a–c
with Fig. 8d–f. The conclusion based on the 20 years
older MORB dataset remains valid and is seen in the
newest data.
We should emphasize that the petrological and geochemical consequences of plate spreading rate variation have long been recognized, such as isotopic
variability (Batiza, 1984), MORB chemical variation
trends (Niu & Batiza, 1993), magma chamber processes
(Sinton & Detrick, 1992), crustal-level thermal and magmatic structures (Chen, 1992; Rubin & Sinton, 2007),
shallow mantle thermal structure (Reid & Jackson,
1981; Niu & Hékinian, 1997a), ridge morphology / topography (e.g. Macdonald, 1982; Macdonald et al., 1998;
Dick et al., 2003) and gravity patterns (Lin & Phipps
Morgan, 1992).
FURTHER INSIGHTS FROM OBSERVATIONS
AND BASIC PETROLOGY
The above has illustrated that to understand the physical controls on ocean ridge processes and MORB
petrogenesis, the logical approach is to isolate the effects of different physical observables. This is conceptually the same as mathematically seeking partial
derivatives for a particular variable while keeping all
other variables held constant (Niu & O’Hara, 2008). For
example, to evaluate the effects of plate spreading rate
variation on MORB petrogenesis, it is best to average
2095
petrological parameters with respect to spreading rate
intervals so that the potential effects of other variables
(e.g. ridge depth variation, mantle source heterogeneity
on varying scales within and between ocean basins)
can be largely or effectively averaged out. Likewise, to
decipher the effects of processes that lead to global
ridge axial depth variation on MORB petrogenesis, it is
best to average petrological parameters with respect to
ridge axial depth intervals so that possible effects by
other variables (e.g. spreading rate variation, dynamic
topography on regional and ridge segment scales as
well as mantle compositional heterogeneity on all
scales within and between ocean basins) can be effectively averaged out. This approach offers us new
insights.
Effects of spreading rate variation on ridge axial
depth and mantle melting
We have demonstrated in Fig. 8a–f that the extent of
mantle melting decreases with decreasing spreading
rate (see Niu & Hékinian, 1997a), which is a straightforward consequence of the ‘lid effect’ as illustrated in Fig.
8g. Because the sub-ridge mantle upwelling rate is, in
an ideal situation, half of the full spreading rate (Phipps
Morgan, 1987; Niu, 1997), the subdued upwelling beneath a slow-spreading ridge should lead to a deep
ridge depth because conductive cooling (1) overcomes
adiabatic upwelling, (2) thickens the cold (and dense)
thermal boundary layer (CTBL), (3) deepens Pf, (4) reduces the PoPf interval, (5) lowers the extent of decompression melting, and predictably (6) produces
thinner magmatic crust. This is correct in terms of basic
petrology and straightforward physics. However, this
anticipated correlation between ridge depth and
spreading rate is not obvious in Fig. 9a. The data points
are still scattered even after plume- or hotspotinfluenced shallow ridges (<1500 m) are removed.
However, as done for MORB chemistry, if we average
ridge depth values with respect to the same spreading
rate intervals, an expected trend emerges as indicated
by the red filled circles.
Figure 9b plots only the spreading rate interval averages with 2s error bars. The topology of the power-law
fitting curve is essentially the same as those of the petrological parameters in Fig. 8, with a flatter trend at
spreading rates >50 mm a–1 and a rapid change
at <40 mm a–1. This is the first-time revelation of the
otherwise long anticipated systematics, and points to
the first-order genetic connections among spreading
rate, ridge depth and petrological parameters.
However, for ridges with spreading rates >50 mm a–1,
the ridge depth is essentially constant, but the petrological parameters define first-order trends that are not
horizontal, and have slopes consistent with a spreading
rate dependent extent of melting (Fig. 8). The average
depth for ridges with spreading rates <40 mm a–1 progressively and significantly deepens with decreasing
spreading rate. Because of the fact that all the average
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Fig. 9. (a) Plot of ridge segment depths as a function of spreading rate using the Gale et al. (2014) (GLD14) data, which show
no simple systematics even when excluding sample sites influenced by mantle plumes or hotspots with depths <1500 m as
indicated (see Fig. 8). Hence, we cannot simply state that ocean
ridge depth increases with decreasing spreading rate. The red
filled circles are ridge depth interval averages as described in
the Fig. 8 caption. (b) As (a), but showing only the depth interval
averages with 2r error bars. The topology of the power-law fitting curve is essentially the same as those of the petrological
parameters in Fig. 8 with a flatter trend at spreading
rates >50 mm a–1 and a rapid change with further spreading
rate decrease. This points to the first-order genetic connections
between spreading rate, ridge depth and petrological parameters. However, for the ridges with spreading rates >50 mm a–1,
the ridge depths define an essentially horizontal band (0 slope;
the purple band), but the petrological parameters define firstorder trends that are not horizontal, and have slopes all consistent with a spreading rate dependent extent of melting (Fig. 8).
The average depths for ridges with spreading rates <40 mm a–1
are progressively and significantly deeper with decreasing
spreading rate (light blue rectangle). Because of the fact that all
the depth averages result from ridge depth interval averages
with the potential effects of all other variables averaged out, the
1100 m depth difference may be the maximum (within uncertainties) ridge depth deepening as the result of reduced spreading rate alone.
ridge depth values are spreading rate interval averages
with the potential effects of all other variables being
averaged out, the 1100 m depth difference may be the
maximum (with uncertainties to be verified) ridge depth
deepening as the result of reduced spreading rate
alone, as indicated in Fig. 9b.
A recent study by Regelous et al. (2016), using an
updated abyssal peridotite dataset and the Gale et al.
Journal of Petrology, 2016, Vol. 57, No. 11&12
(2014) MORB dataset, confirmed the recognition by Niu
& Hékinian (1997a; also Niu & O’Hara, 2014) that the extent of ridge mantle melting decreases with decreasing
spreading rate, especially beneath slow-spreading
ridges as demonstrated above (Fig. 8). Regelous et al.
further concluded, however, that this rules out the effect
of fertile mantle compositional variation. This latter
statement is imprecise and incorrect. Let us examine
why this statement needs correction. In Fig. 7 the correlations of MORB chemistry with ridge depth are best interpreted as mantle major element compositional
variation as per Niu & O’Hara (2008), and provide no information on spreading rate effect because this effect is
averaged out in these plots. The Fig. 8 correlations of
MORB chemistry with spreading rate are consistent
with spreading rate-dependent varying extent of melting (as shown by these authors), but yield no information on ridge depth effects because these are again
averaged out. Hence, on the basis of Fig. 8 we cannot
rule out the effect of ridge depth and mantle source effects. Now, we can see by comparing Fig. 8 and Fig. 9
that the largest MORB compositional variation occurs at
slow spreading rates of <40 mm a–1, which corresponds, conservatively, to a ridge axial depth variation
of 3200 m to 4200 m. Within this depth range, we see
no constant, but a significantly systematic MORB compositional variation in Fig. 7.
Therefore, it is conceptually important and correct to
state that globally primitive MORB major element compositions (i.e. MORB melts with Mg# > 072) are largely
controlled by both fertile mantle compositional variation (e.g. Fig. 7) and plate spreading rate variation
(Fig. 8) with the spreading rate effect being especially
prominent at very slow-spreading ridges (Figs 8 and 9).
MORB melts preserve no signature of Po and To,
but signature of Pf and Tf
By stating that global MORB major element compositional variations (e.g. Fe8, Fe90) result from mantle (potential) temperature variations of 250 K (Klein &
Langmuir, 1987; Langmuir et al., 1992; Dalton et al.,
2014; Gale et al., 2014), this specifically refers to variation of the depth at which upwelling mantle intersects
the solidus; that is, the initial depth of melting Po and
the corresponding To, as schematically shown in
Fig. 10. Prior to the introduction of the final depth of
melting Pf and the corresponding Tf (Niu & Batiza,
1991a), the ridge mantle decompression melting was
thought to continue up to the base of the crust (Klein &
Langmuir, 1987; McKenzie & Bickle, 1988). It follows
that a hotter parcel of upwelling mantle would intersect
the solidus deeper, have a longer melting column and
produce more melt with the signature of higher pressure (high Fe8) and higher extent (low Na8) of melting,
whereas a cooler parcel of upwelling mantle would
intersect the solidus shallower, have a shorter melting
column and produce less melt with the signature of a
Journal of Petrology, 2016, Vol. 57, No. 11&12
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Fig. 10. Schematic illustration (left) and qualitative illustration (right) of sub-ridge thermal structure with explanation of the elements (see Figs 7 and 8). The plate separation induced decompression melting begins when the upwelling mantle intersects the solidus at Po (at corresponding To) and continues until the upwelling mantle reaches Pf (at corresponding Tf), the final depth of melting
or melt–solid equilibration, which is the base of the cold thermal boundary layer beneath the ridge. Because of its buoyancy, the
melt (red-arrowed thin dashed lines) ascends faster than the residual solids (blue-arrowed thick lines). The melt is extracted to form
the ocean crust whereas the residue contributes to the lithospheric mantle beneath the crust (shallow portions sampled as abyssal
peridotites). The globally large mantle (potential) temperature variation of 250 K claimed by Klein & Langmuir (1987) and Gale
et al. (2014) to be preserved in MORB chemistry refers to large To and Po variation. The ‘lid effect’ argued by Niu & O’Hara (2008) to
be preserved in MORB chemistry here refers to Tf and Pf, either because of mantle density and related decompression melting (Fig.
7; Niu & O’Hara, 2008) or because of plate spreading rate variations and related decompression melting (Fig. 7; Niu & Hékinian,
1997a,b). It should be noted that because of melt–solid re-equilibration during decompression melting (PoPf) on a time scale of
thousands of years (re-equilibration is achieved in tens of hours experimentally), the MORB melts do not record the signature of Po
and To; hence there is no petrological evidence for large mantle potential temperature variation of 250 K beneath global ocean
ridges.
lower pressure (low Fe8) and lower extent (high Na8) of
melting.
Niu and co-authors (Niu, 1997, 2004; Niu & Hékinian,
1997a,b; Niu & O’Hara, 2008) have emphasized that
there is no evidence for large mantle potential temperature variation beneath global ocean ridges. The petrological parameter Fe8 (now also Fe90) as a pressure
indicator collapses (see above). Dalton et al. (2014) used
ridge axial depth variation as evidence for large mantle
potential temperature variation, but this is a circular argument and has no foundation (see above and below)
because fertile mantle density variations owing to compositional variations (compositional buoyancy variation) are much more effective in causing ridge axial
depth variation. Niu and co-authors advocate the presence and significance of the cold thermal boundary
layer (CTBL) beneath ocean ridges because this is consistent with physical analysis (Reid & Jackson, 1981)
and petrological observations (Niu, 1997, 2004; Niu &
Hékinian, 1997a,b; Niu et al., 1997; Niu & O’Hara, 2008)
as elaborated above (Figs 7 and 8). The thickness of the
CTBL determines the final depth of melting or melt–
solid re-equilibration, Pf, hence affecting the decompression melting interval PoPf and the extent of melting. This is called the ‘lid effect’ (Niu et al., 2011) as
elaborated above as functions of fertile mantle compositional / density variation and plate spreading rate variation with correlated variations of MORB chemistry
with ridge axial depth (Fig. 7) and plate spreading rate
(Fig. 8).
The key message of Fig. 10 concerns the very fundamental question as to whether the erupted MORB melts
can preserve the signature of Po and To. The answer is
simply ‘No’. The melt formed by decompression (in the
depth interval PoPf) will ascend because of buoyancy
and melt segregation caused by matrix compaction
(McKenzie, 1984). This occurs as soon as the initial porosity exceeds the permeability threshold, the value of
which has varied from up to 10% (e.g. Maaloe, 1982), to
3% (McKenzie, 1984), and to several orders of magnitude less than 1% (McKenzie, 1985; Spiegelman &
Elliott, 1993) to explain the observed Th–Ra excess in
MORB. However 1% porosity may be more consistent
with physical analysis (Turcotte & Phipps Morgan,
1992) and petrological constraints (Johnson et al.,
1990). Melt migration and ascent may occur as diffuse
porous flow or network-like channeling on millimeter
scales (e.g. McKenzie, 1984; Turcotte & Phipps Morgan,
1992).
Therefore, recognizing that (1) porous flow is the primary means of melt transport, (2) the time scale for
melt transport from the melting region to the crust is of
the order of thousands of years (e.g. McKenzie, 1984,
1985; Spiegelman & McKenzie, 1987; Rubin &
Macdougall, 1988, 1990), and (3) a time of no more than
tens of hours is sufficient for attaining solid–melt
2098
equilibrium in peridotite melting experiments, lowpressure melt–solid equilibration is inevitable during
melt ascent (Niu, 1997). Hence, the erupted melts will
not preserve the signature of the initial depth of melting
(Po, To), nor the signatures of any condition in the decompression interval PoPf. The pressure signature the
erupted MORB melts can retain would be the final depth
of melting or melt–solid equilibration of Pf and Tf, which
is the condition at the base of the CTBL (¼ ‘lid’). The
MORB compositional variation as a function of ridge
depth variation (composition-controlled density variation; Fig. 7) and plate spreading rate variation (Fig. 8)
is the manifestation of the ‘lid effect’; that is, the Pf and
Tf signature preserved in MORB melts (Figs 7, 8 and 10).
Because the erupted MORB melts have lost the signature of the initial depth of mantle melting (i.e. Po and
To) owing to the inevitable melt–solid re-equilibration,
how can we use MORB chemistry to say that large mantle potential temperature variations (250 K) exist beneath global ocean ridges? How can we logically
establish primary magmas parental to MORB with MgO
as high as 19 wt % (corrected to Fe90; see Fig. 3d)? In
Fig. 10 it is shown explicitly that the melt produced will
ascend in the form of porous flow (< 2%) through the
solid peridotite matrix (>98%) and undergo continuous melt–solid re-equilibration. The re-equilibration is
expected to be complete during decompression melting
(PoPf) because of the close melt–solid contact and because the thousands of years of melting and melt segregation is long enough compared with the tens of hours
required to reach equilibrium in experimental charges,
especially for elements such as Si, Fe and Mg controlled
by olivine, which is the dominant mantle peridotite mineral. Niu (1997) presented this argument explicitly, but I
ask readers to judge objectively whether MORB melts
can preserve the signature of Po and To in terms of Si,
Fe and Mg as presumed by Klein & Langmuir (1987),
Langmuir et al. (1992), Dalton et al. (2014) and Gale
et al. (2014).
A point that should be remembered: the
meaning of the correlated compositional
variations of MORB and abyssal peridotites
We should note the nature of the compositional correlation between abyssal peridotites and spatially associated MORB (Dick et al., 1984) and the petrological
interpretations of this correlation (Dick et al., 1983; Klein
& Langmuir, 1987). This correlation was effectively
summarized in fig. 1 of Niu et al. (1997), showing a scattered but statistically significant positive linear correlation between the modal clinopyroxene (vol. %) content
of abyssal peridotites and the Na8 of spatially associated MORB (18 site averages with R2 ¼ 074). Niu et al.
(1997) proposed that when sample locations approach
‘hotspots’ with shallowed ridge depths, both Na8 and
modal clinopyroxene decrease, which is consistent with
an increasing extent of mantle melting, influenced by
hotter mantle beneath hotspots. This observation is
Journal of Petrology, 2016, Vol. 57, No. 11&12
simple and powerful. The interpretation is logical also,
but may not be correct. Following a detailed bulk-rock
major and trace element study of global abyssal peridotites, Niu (2004, p. 2452) could not avoid the conclusion
that: ‘there is no clear justification that Na8 in MORB or
cpx mode in abyssal peridotites genuinely reflects the
extent of mantle melting beneath mid-ocean ridges, nor
that such an inferred extent of melting reflects subridge mantle potential temperature variations. The complementarity could very well be inherited from the fertile mantle sources.’ This was echoed by Dick et al.
(2007), who stated: ‘As the previously analyzed peridotites are from fracture zone walls 05 to 14 Ma. old, and
the “spatially associated basalts” are largely from the
modern ridge axis, this argues for a long-term stability
in magma composition and therefore mantle composition as well.’ Indeed, despite the compositional complementarity between MORB (melts) and abyssal
peridotites (residues), it is unlikely for the 14 Myr old
abyssal peridotites to be melting residues of the
present-day MORB. Mantle thermal anomalies may be
long-lived, but can decay away with time. However,
compositional anomalies (and buoyancy) should be
long-lasting. For this, Zhou & Dick (2013) provided a
good case study.
The geodynamic significance of compositional buoyancy contrast is best manifested by the largest topographic contrast on the Earth: the compositionally
buoyant continental lithosphere (the crust and mantle
lithosphere) and the compositionally dense oceanic
lithosphere (despite the contribution of the ‘coldness’ of
the latter) (Niu et al., 2003; Niu, 2014, 2016).
SUMMARY
1. The popular petrological parameter Fe8 devised by
Klein & Langmuir (1987) and advocated by
Langmuir et al. (1992) and Gale et al. (2014) in
studying MORB mantle melting conditions is misleading, as demonstrated by Niu & O’Hara (2008),
because Fe8 represents variably evolved melts with
Mg# ¼ 056–068 (Figs 1–3), which is a crustal signature and cannot be in equilibrium with mantle olivine. Recent work by Gale et al. (2014) showed that
Fe8 Fe90 (Fig. 3b), with the purpose of advocating
Fe8 to be the same as Fe90, and to reflect melts in
equilibrium with mantle olivine Fo90. This is not
true. Because melts with Fe8 have Mg# ¼ 056–068,
but those with Fe90 are meant to have Mg# 072–
073, how can Fe8 Fe90? (Fig. 3b). The data treatment to make Fe90 ! Fe8 was achieved by adding
varying amount of olivine, but can MORB melts
have MgO as high as 19 wt % (Fig. 3d)? The most
primitive MORB melts with Mg# 072 have
105 wt % MgO (data constrained). The Gale et al.
(2014) corrected Fe90 (¼ Fe8) values have no relevance to actual data as they are compositionally
remotely different from any MORB composition
Journal of Petrology, 2016, Vol. 57, No. 11&12
and are beyond petrologically permitted interpolation and extrapolation (Fig. 3d).
2. Global MORB major element compositional systematics in MgO or Mg# variation diagrams are not
simple liquid lines of descent (LLD; Figs 4 and 5),
but the net effect of cooling-dominated, crustallevel processes (e.g. fractional crystallization,
magma mixing, melt–rock assimilation / reaction
and other aspects of complex open-magma chamber processes). It is petrologically justified to correct MORB melts for the crustal-level effects to
Mg# 072 to be in equilibrium with mantle olivine
of Fo 90 by using a single set of correction coefficients applicable to the global MORB dataset.
Global MORB FeO values corrected to Mg# ¼ 072
have Fe72 7–8 wt % (i.e. 10 Fe units), which is
well constrained by the global MORB data as illustrated in FeO–Mg# diagrams (Figs 4 and 5). This
new demonstration using the actual data as done
by Niu & O’Hara (2008), argues against the incorrect statement by Gale et al. (2014) that the small
Fe72 range is due to an inaccurate correction procedure. In contrast, regardless of whether the correction procedure by Gale et al. (2014) to Fe8 is
justified / accurate or not, the new global MORB
dataset shows that Fe8 (¼ FeO at MgO ¼ 80 wt %) is
7–125 wt % (> 55 Fe units), corresponding to
Mg# ¼ 056–069 (Figs 3–5), which cannot be used
to say anything petrologically meaningful about
mantle processes.
3. The correlation between MORB chemistry (Ti72,
Al72, Fe72, Mg72, Ca72, Na72, Ca72/Al72) and ridge
axial depth (Fig. 7) indicates a genetic relationship
between the two, which is not a ‘cause-and-effect’
one; instead, both are different effects of a common
cause. The common cause is sub-ridge fertile mantle compositional variation, which, by inference, is
most consistent with varying extent of prior melting and melt extraction in terms of major elements.
The least depleted (or relatively enriched) MORB
source is characterized by higher TiO2, Al2O3 and
Na2O, and lower FeO, MgO and CaO, and thus
lower CaO/Al2O3 than the most depleted MORB
source. Such major element compositional variation determines the mantle mineralogy with
greater bulk-rock density for the progressively least
depleted (or relatively enriched) source, leading to
progressively deeper ridges because of isostasy
(Fig. 7). Melting of such compositionally varying
sources imparts the source chemical signature to
the erupted MORB melts (Fig. 7a and b), leading to
the observed correlation of MORB chemistry with
ridge axial depth (Fig. 7).
4. Mantle density is an important physical property
and its variation must have a varying control on
plate separation-induced upwelling. With increasing sub-ridge mantle density, as expressed by the
increasing ridge axial depth, the maximum extent
of passive upwelling is progressively subdued,
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5.
6.
7.
8.
9.
resulting in a slightly reduced decompression melting interval and somewhat lower extent of melting.
This has an enhanced (positively superimposed) effect on the signature of source compositioninherited MORB chemistry and the correlation with
ridge axial depth (Fig. 7).
The updated global MORB dataset provided by
Gale et al. (2014) confirms the recognition 20
years earlier by Niu & Hékinian (1997a) that the extent of sub-ridge mantle melting increases with
increasing plate spreading rate, which is expected
in terms of the understood physics of the way in
which ocean ridges work (Fig. 8), but disproves the
conclusive statement by Gale et al. (2014): ‘There is
no correlation between the chemical parameters
and spreading rate.’
Mantle temperature variation could play a role in
MORB petrogenesis, but its overstated role comes
from the misleading parameter Fe8 (and Fe90), from
using ridge axial depth as a constraint by neglecting the intrinsic chemical and physical controls of
fertile mantle compositional variation, and from
treating ridge-centered mantle plumes such as the
Iceland plume as being of plate tectonic origin.
In addition to the incorrect data treatment by Gale
et al. (2014), the argument for large mantle temperature variations has an erroneous presumption;
that is, MORB melts retain the signature of their initial depth of mantle melting (Po and To) in terms of
Fe8 (and Fe90). The decompression melting begins
at Po and continues until Pf, during which the melt
phase is continuously formed (<2%) and tends to
ascend because of buoyancy, while re-equilibrating
with the peridotite matrix (>98%) on a time scale of
thousands of years. The continued re-equilibration
is inevitable because of the long enough time for
close physical contact between the melt and solid,
as is readily achieved experimentally in tens of
hours. This is particularly true for Si, Mg and Fe
controlled by olivine, the dominant phase of the
peridotite matrix (Fig. 10). Therefore, MORB melts
can only retain the final depth of melting or melt–
solid equilibration (i.e. Pf and Tf), not To and To, as
explicitly discussed previously (Niu, 1997) and
demonstrated by the fact that the most primitive
MORB melts with Mg# 072 have MgO 105 wt
%, not higher and not as high as 19 wt % (Fig. 3d).
The preservation of Pf and Tf in MORB melts illustrates that the lid effect, which has a primary control on the extent of melting, final depth of melt–
solid equilibration and melt compositions for intraplate ocean island magmatism (Humphreys & Niu,
2009; Niu et al., 2011), also exerts a key control on
MORB petrogenesis as manifested by the correlations of MORB chemistry with ridge axial depth
(Fig. 7c) and with ridge spreading rate (Fig. 8g).
The correlations of MORB chemistry with ridge
axial depth cannot be used to rule out a plate
spreading rate control. Likewise, the correlations of
2100
MORB chemistry with plate spreading rate variation cannot be used to rule out a fertile mantle
compositional control. Both fertile mantle compositional variation (expressed as large ridge axial
depth variations owing to composition-determined
density variations) and spreading rate variation are
two primary variables that control first-order ocean
ridge processes in general and MORB petrogenesis
in particular.
10. It has long been anticipated that ridge axial depth
should increase with decreasing plate spreading
rate, but the lack of correlation between the two
has been puzzling. This correlation does indeed
exist after removing hotspot-influenced ridges and
after averaging ridge depth values within spreading
rate intervals (Fig. 9). This correlation is best
described by a non-linear power-law relationship
that also best describes the relationship between
MORB chemistry and spreading rate. All these observations are straightforward consequences of
plate tectonics.
11. The meaning of the global MORB major element
compositions may continue to be debated, but the
conclusions offered here are the most objective
and logical, and are most consistent with petrological, geochemical, geological and geophysical principles and observations.
12. To promote scientific debate, I end this paper here
by sharing my philosophy: for the same data, there
can be many different interpretations, but reasonable interpretations are limited and, strictly speaking, the correct interpretation must be openminded and objective.
FUNDING
This work was supported by Chinese NSF (Grants
41630968 and 91014003), Chinese NSF–Shandong
(Grant U1606401), Chinese Academy of Sciences
(Innovation Grant Y42217101L), Qingdao National
Laboratory for Marine Science and Technology (Grant
2015ASKJ01), and grants from regional and local
authorities (Shandong Province and City of Qingdao).
ACKNOWLEDGEMENTS
I write this paper in commemoration of Mike O’Hara for
his friendship, for our scientific exchanges and for the
memorable times we had over the previous 18 years
before his passing through emails, phone calls, lunchtime chats and extended visits. His constructive review
of my abyssal peridotite paper for this journal (Niu,
1997) began our productive research collaboration and
led to my migration from the University of Queensland
to Cardiff and to Durham. We published 10 papers together on topics including subduction initiation, continental crust accretion, petrogenesis of intra-plate ocean
island basalts, ocean ridge basalts and lunar rocks. Our
friendly discussion and scientific debates benefited me
Journal of Petrology, 2016, Vol. 57, No. 11&12
tremendously. It may surprise some, but we actually
disagreed more often than agreed and he was extremely happy when he was convinced that I was correct. An anonymous journal reviewer is thanked for the
very detailed and thorough constructive comments on
an early version of the paper. The comments and editorial efforts by Marjorie Wilson and Alastair Lumsden
have improved the paper, for which, and for their exceptional patience, I am grateful.
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